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Storage conditions and eruptive dynamics of central
versus flank eruptions in volcanic islands: the case of
Tenerife (Canary Islands, Spain)
Joan Andújar, Fidel Costa, Bruno Scaillet
To cite this version:
1
Storage conditions and eruptive dynamics of central versus flank
eruptions in volcanic islands: the case of Tenerife (Canary Islands,
Spain)
Joan Andújara,*, Fidel Costab, Bruno Scailleta
a.
Université d’Orléans, ISTO, UMR 7327, 45071, Orléans, France ; CNRS/INSU, ISTO, UMR 7327, 45071 Orléans, France ; BRGM, ISTO, UMR 7327, BP 36009, 45060 Orléans, France.b
. Earth Observatory of Singapore, Nanyang Technological University, Singapore 639798, Singapore* Corresponding author: Joan Andújar.
phone number: (+33) 2 38 25 53 87
Fax: (+33) 02 38 63 64 88
e-mail address: Juan.Andujar@cnrs-orleans.fr
Fidel Costa e-mail address: fcosta@ntu.edu.sg
Bruno Scaillet e-mail address: bscaille@cnrs-orleans.fr
KEY WORDS: Phase equilibria, phonolite, experimental petrology, eruptive dynamic,
Abstract
We report the results of phase equilibrium experiments on a phonolite produced during one of the most voluminous flank eruptions (ca. 1 km3) of the Teide-Pico Viejo complex (Tenerife Island). Combined with previous experimental and volcanological data we address the factors that control the structure of the phonolitic plumbing system of Teide-Pico Viejo stratovolcanoes. The Roques Blancos phonolite erupted ca 1800 BP and contains 14 wt % phenocrysts, mainly anorthoclase, biotite, magnetite, diopside and lesser amounts of ilmenite. Crystallization experiments were performed at temperatures of 900ºC, 850ºC and 800ºC, in the pressure range 200 MPa to 50 MPa. The oxygen fugacity (fO2) was varied between NNO+0.3 (0.3 log units above to the Ni-NiO solid
buffer) to NNO-2, whilst dissolved water contents varied from 7 wt% to 1.5 wt%. The comparison between natural and experimental phase proportions and compositions, including glass, indicates that the phonolite magma was stored prior to eruption at 900±15ºC, 50±15MPa, with about 2.2 wt% H2O dissolved in the melt, at an oxygen
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1. Introduction
Understanding the plumbing system beneath volcanic edifices is crucial to constrain the parameters that control the evolution of ascending magmas and for anticipating the future eruptive behaviour of the volcano. The use of dense geophysical and geochemical monitoring networks deployed on highly active volcanes (i.e., Etna volcano; Bonacorso et al. 2004) allows to gain information concerning the movement and likely levels of magma storage. Monitoring data combined with petrological studies (e.g., Kahl et al., 2011) enables to better reconstruct the geometry of the plumbing system beneath the volcanic edifice of an on-going eruption, and thus construct more robust eruptive scenarios in the short time frame.
However, real time monitoring techniques do not easily allow to understand the mid (e.g., hundred years) to long-term (thousands of year or more) evolution of the volcanic system, or the variation of eruptive dynamics (ie.; variation in eruptive style during a single event or between eruptions) with time. This is particularly true for volcanoes with low eruption frequency, which are characterised by long periods of dormancy during which little, if any, geophysical or geochemical signals can be recorded, making their interpretations difficult whenever the system awakes. The volcanological record often bears witness of significant variations in the eruptive styles (alternation between sustained explosive Plinian eruptions, transient explosive and purely effusive activity), and in vent location (central and flank eruptions; i.e., Corsaro et al., 2007; Ablay and Martí, 2000), which may in part reflect the complexity of the plumbing system. In such cases, only the study of past eruptions and the distribution of eruptive centres can provide insights on both the long-term evolution of the volcanic edifices and on the factors that control the eruptive dynamic.
Rodríguez-Badiola, 2006; Carracedo et al., 2007; Martí et al., 2008; Wiesmaier et al., 2012). Recent numerical simulations performed on this volcanic complex suggest that the main factors controlling the vent location at Teide-Pico Viejo are the shape of the reservoir and the stress field distribution around it which can be affected by the presence of a second magma chamber (Martí and Geyer, 2009). Such a possibility is supported by the petrological and geochronological data of recent phonolitic eruptions (Ablay et al., 1998; Carracedo et al., 2007; Andújar and Scaillet, 2012a). The relative position between the reservoirs also exerts an important control on the final trajectory of the injected dykes and hence in the final location (central or flank) of the vent (Martí and Geyer, 2009). Since the depths of such reservoirs feeding flank eruptions at Teide-Pico Viejo are poorly constrained (Andújar and Scaillet, 2012a), we have performed phase equilibrium experiments on a representative phonolite of the Roques Blancos, one of the most recent flank dome eruptions occurred at this volcanic island.
2. Geological setting
The evolution and stratigraphy of Teide stratovolcano have been the focus of several studies (i.e., Araña, 1971; Martí et al., 1994; Ablay et al., 1998; Martí and Gudmunsson, 2000; Ablay and Martí, 2000; Rodríguez-Badiola et al., 2006; Carracedo et al., 2007; Wiesmaier et al., 2012). We first summarise the main aspects concerning the structural and volcanological evolution of the actual volcanic complex of Teide and Pico Viejo stratovolcanoes.
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either at central Pico Teide or from the numerous flank vents (up to 10), of which the Montaña Blanca, Roques Blancos, Pico Cabras and Montaña Rajada events are the most important ones (Fig.1b; Ablay and Martí, 2000; Rodríguez-Badiola et al., 2006; Carracedo et al., 2007).
Although such satellite eruptions are volumetrically smaller (generally ≤ 0.2 km3) relative to those from the central Teide vent (generally ≥ 0.3 km3), they are characterised by a sustained explosive activity and fall-out deposits together with the emission of thick lava flows (García et al., 2012; Martí et al., 2012). The explosive activity of these flank eruptions includes the sub-Plinian event of Montaña Blanca (Ablay et al., 1995; Ablay and Martí, 2000; Rodríguez-Badiola et al., 2006; Martí et al., 2008; Andújar and Scaillet, 2012a), and El Boqueron flank eruption (García et al., 2012). In addition to these episodes, two more phonolitic explosive flank events have been recently identified (Martí et al., 2012; O. García unpublished data). In contrast, phonolitic volcanism from Teide proper is less explosive, and alternates between pure effusive activity that generated thick lava flows and transient explosive activity, which has produced scoria and spatter deposits (Ablay and Martí, 2000; Martí et al., 2008). However, it is unclear why some eruptions proceed through the central Teide cone while others occur on its flanks, and how does this relate to the more explosive activity from the flank events. Here we provide petrological constraints to explain these observations by determining the pre-eruptive conditions of the Roques Blancos dome, one of the most significant flank eruptions in the recent history of Teide-Pico Viejo volcanic complex, in terms of the erupted volume and extension of the lava flows (Fig.1; Balcells and Hernández-Pacheco, 1989; Carracedo and Rodríguez-Badiola, 2006; Martí et al, 2008).
3. The Roques Blancos eruption
part of Tenerife Island (Carracedo and Rodríguez-Badiola 2006; Martí et al., 2012). Both phonolitic products are petrologically and geochemically very similar (Fig.1; Balcells and Hernández-Pacheco, 1989; Ablay, 1997; Ablay et al., 1998; Ablay and Martí, 2000; Carracedo and Rodríguez-Badiola, 2006; Carracedo et al., 2007; Martí et al., 2008).
3.1. Pre-eruptive conditions of Roques Blancos magma from petrological observations
Several blocks of lava were collected at 2500 m asl from one of the main lava flows of the 1714 BP Roques Blancos youngest dome (Fig.1). The freshest rock blocks were selected as starting material for performing phase equilibrium experiments, whole-rock, and electron microprobe analyses (EMPA) of the different phases (Table 1). The bulk-rock composition, as determined by ICP-MS, is a typical phonolite (59.4 wt% SiO2, 15.4 Na2O+K2O; Le Maitre et al., 1989; Table 1) representative of the Roques
Blancos flows (Balcells and Hernández-Pacheco, 1989; Ablay et al., 1998; Carracedo and Rodríguez-Badiola, 2006).
Modal point counting of three thin sections (3000 points each) was combined with mineral densities to obtain the weight percent (wt%) of the different phases. The sample has 14 wt % phenocrysts, mainly anorthoclase (13.7 wt%; An4-3, Ab69-65, Or 32-27), minor amounts of biotite (0.3 wt%; Mg#62-63; Mg#=100[Mg/(Mg+Fe*)]),
magnetite (0.3 wt%; Mg# 5), diopside (0.1 wt%; Mg# 68-70; En39-41,Fs9-11,Wo50) and
ilmenite (0.1 wt%). Phenocrysts are set in a highly crystalline groundmass made of microlites of alkali feldspar, magnetite, clinopyroxene, in addition to glass (Table 1, Fig.2). Mass-balance calculations were used to obtain the composition of the residual melt that was in equilibrium with phenocrysts by combining the whole-rock and mineral phase compositions. As anticipated owing to the low phenocryst content, the resulting melt is phonolitic (59.2 wt% SiO2, 17.0 wt% Na2O+K2O) with a composition similar to
that of the bulk-rock (Table 1).
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calcium-and Ba-rich rims of An7-8 and up to 0.4 wt%, respectively; Fig.3). In detail,
many crystals show complex rims: (1) the outermost rims (e.g., last 10-20 m) have extremely high Fe and Ti, high Ca, and very low Ba, possibly resulting from fast growth, during which incompatible elements are incorporated at higher than equilibrium concentrations, and compatible ones at lower than equilibrium concentrations (e.g., Albarède and Bottinga, 1972). (2) Inside this outer rim, crystals still have a high Fe, Ti, Ca and Ba zones, compared to the core. There is a dissolution zone between the inner part and the outer rims of the crystals. These observations suggests that there was a change in the pre-eruptive conditions just prior to eruption, ie magma reheating upon a recharge event, followed by a fast crystal growth stage.
Pre-eruptive temperatures and fO2 for the 1714 BP Roques Blancos phonolite
were determined by using six co-existing Fe-Ti oxides and the thermo-oxybarometric model of Sauerzapf et al. (2008). Results yield temperatures of 895±5°C and a log fO2
of -12.2, equivalent to an oxygen fugacity of 0.2 log units below the Ni-NiO solid buffer (NNO-0.2; Table 1). As there are currently no constraints on storage pressure or volatile content of the Roques Blancos phonolite, the ranges of pressure and volatile content covered in the experiments were guided by previous results obtained on similar compositions of the same volcanic complex (Andújar et al., 2010; Andújar and Scaillet, 2012a).
4. Experimental work
We have performed phase-equilibrium experiments following a methodology similar to that used for constraining the storage conditions of active volcanoes such as Stromboli, Vesuvius, St. Pedro volcanoes (e.g., Di Carlo et al., 2006; Scaillet et al., 2008; Costa et al., 2004) and taking in account the recommendations in the application of phase equilibrium experiments to volcanic rocks as laid down by Pichavant et al. (2007).
4.1. Preparation of the starting material
significant Na or Fe loss compared to the starting rock (Table 1). The dry glass was then ground for obtaining a powder that was used as starting material for hydrothermal experiments and stored in an oven at 120ºC.
4.2. Experimental equipment and procedures
A total of 55 crystallization experiments were performed at the experimental laboratory of ISTO (Orléans, France) in a vertically operating Internally Heated Pressure Vessel (IHPV), using Ar as pressurising medium which was mixed with different amounts of H2 at room temperature in order to achieve different fO2 (Scaillet et
al., 1992). A transducer calibrated against a Heise Bourdon gauge with an uncertainty of ±2 MPa was used for recording total pressure. Experiments were performed using double-winding molybdenum and kanthal furnaces which produce near-isothermal conditions (gradient <2-3ºC/cm) along a 3 cm long hot spot. S- or K-type thermocouples with an accuracy of ±5ºC were used to record experimental temperature and a rapid-quench technique was systematically used which allows isobaric quenches at high cooling rates (>100ºC/s) (e.g., Martel et al., 1999; Di Carlo et al., 2006). In all runs reported here, the drop quench was successful as indicated by the rise in total pressure upon the falling of the sample holder into the cold (bottom) part of the vessel.
Experiments were mainly conducted at 850 ºC and 900 ºC, pressures of 200, 150, 100, and 50 MPa and at fO2 NNO+0.3 to NNO-2. The effect of temperature on
phase relationships and compositions was studied by conducting experiments at 800ºC, 200 and 100 MPa at NNO. One run was also performed at NNO-1 at 850 ºC and 100 MPa so as to study the effect of oxygen fugacity on phase equilibria at this temperature (Table 2).
4.2.1. Capsule preparation
H2O-saturated and undersaturated charges were prepared using 1.5 cm long, 2.5 mm
inner diameter, 0.2 mm wall thickness Au capsules. Distilled H2O was first loaded, then
silver oxalate as the source of CO2 for H2O-undersaturated charges, and finally the glass
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have explored various starting H2O/CO2 ratios at a given T-P conditions: XH2Oin,
defined as the H2O/(H2O+CO2) molar ratio, was varied in the range 1-0.48 (Table 2).
For each run, one capsule containing a Ni-Pd-O alloy sensor, prepared following the procedure of Taylor et al. (1992), was added to monitor the fH2 during the experiment. A
typical run consisted of loading the vessel with six capsules (each with the same starting material but different XH2Oin) and the Ni-Pd-O sensor capsule, all of them experiencing
the same T-P-fH2 conditions (Table 2).
Depending on pressure and temperature, the run duration varied between 7 and 18 days (Table 2). Runs were terminated by first using the drop quench device and then switching off the power supply. After the experiment, capsules were checked for leaks, opened, and half of the run product was mounted with epoxy resin and polished for optical observation, and subsequent EMP and SEM analyses. The same procedure was followed for the capsules containing the metallic sensors and the subsequent EMP analysis of the metallic pellet allowed the determination of the fO2 of the system
(Pownceby and O’Neill, 1994; see below).
4.2.2. Water content, fH2, fO2 in the capsules
The use of different mixtures of H2O+CO2 in the capsules allowed us to explore
different water fugacities, hence, different melt H2O content in the experiments (Table
2). In experiments conducted at 900ºC and at H2O-saturated conditions, the size of glass
pools was large enough for determining the amount of dissolved water by Fourier transform infra-red (FTIR) spectroscopy. The composition of the natural phonolite resembles that used by Carroll and Blank (1997) for determining the solubility of water in phonolitic melts. Thus, we have used the same analytical conditions and parameters (e.g., extinction coefficients 5200 4500) to determine the water concentration in the glass (Table 3). The small size of the glass pools in experiments run at 850-800ºC did not allow for water determination using FTIR. In such cases, we have calculated the water content of the experimental charges by using the H2O solubility model of Papale
et al. (2006).
Once the water content of the glass in H2O-saturated charges is known, the water
content of charges held at the same temperature and pressure but at XH2Oin<1, was
calculated by multiplying the water content determined at H2O-saturated conditions by
We used the Ni-Pd-O sensors from successful runs to determine the prevailing
fH2. The fO2 was then determined by using the dissociation constant of water (Robie et
al., 1979) and knowing the fH2O at the experimental temperature and pressure. The
fH2O was calculated as fH2O=XH2Oin*fH2Oo where fH2Oo is the fugacity of pure water
(in bars) at the running temperature and pressure (Burnham et al., 1969). In experiments where the sensors failed, the fH2 was determined using an empirical calibration curve
between the H2 pressure added to the autoclave at room temperature and the fH2
retrieved from runs in which the sensor was successful (Table 2).
The experiments were conducted at oxygen fugacities around NNO±0.5. However, as fO2 varies with decreasing activity of water in the melt [aH2O, (or XH2Oin
in the ideal approximation and neglecting the amount of water dissolved in the melt)] at fixed T, P, fH2 (e.g., Scaillet et al., 1995; Freise et al., 2009; Andújar and Scaillet,
2012a), water-undersaturated experiments were effectively run at somewhat lower fO2
(NNO-1±0.5 log units ; Table 2). We have also conducted experiments at 850ºC and 100 MPa and at an fO2 1 log unit lower than the average conditions so as to determine the
effects of fO2 on phase relationships.
4.2.3. Analytical techniques
Phase compositions were determined using a Cameca SX-50 electron microprobe with an accelerating voltage of 15 kV, sample current of 6 nA, and a counting time of 10 s. A defocused beam of 10 µm was used for glasses and a focused beam for minerals. Alkali migration was corrected by using secondary phonolitic standards with a composition similar to the natural obsidian and with known dissolved water contents of 10 wt%, 6 wt% and 1.5 wt% respectively (Andújar et al., 2008, 2010). Mass-balance calculations were used to obtain the phase proportions of the charges using the bulk composition of the starting material and the composition the phases (Table 2).
Experiments conducted at 900ºC and 850ºC had small crystals (< 10-5 µm) which did not allow to obtain EMP analyses free of glass contamination (i.e.; clinopyroxene with K2O higher than 0.1 wt%, see below). The glass contribution was
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4.2.4. Attainment of equilibrium
The crystallization experiments show textural and compositional features similar to those observed in crystallization experiments conducted on phonolitic compositions (e.g., Freise et al., 2003; Scaillet et al., 2008; Andújar et al., 2008, 2010; Fig. 2b). These features include: (1) a homogeneous distribution of the phases within the charges, (2) euhedral crystals, (3) homogeneous phase compositions, including glass (Fig. 2), and (4) the smooth variation of phase proportions and compositions with changes in experimental conditions. The duration of our experiments (1 to 2 weeks) is within the range of that applied in other studies, including for phonolitic systems (Berndt et al., 2001; Freise et al., 2003; Harms et al., 2004; Scaillet et al., 2008), for which close to equilibrium conditions were also proposed (see also Scaillet and Evans 1999; Costa et al., 2004; Andújar et al., 2008, 2010, and Pichavant et al., 2007).
5. Experimental results
5.1. Phase relations
The main features of the phase relations are shown in a series of polybaric-isothermal (Fig. 4a-c) or isobaric-polythermal (Fig. 4d) sections. At 900ºC the liquidus phase is biotite at all investigated pressures and melt water content (H2Omelt). Biotite
crystallization is followed by the co-crystallization of magnetite plus ilmenite, then alkali feldspar with decreasing H2Omelt. Clinopyroxene has a narrow stability field and
only appears at H2Omelt < 2.5 wt% at 900 ºC. At about 850ºC, magnetite and ilmenite
become the liquidus phases and biotite co-crystallization occurs at H2Omelt < 6 wt%
(Fig. 4). At 850 ºC the stabilities of alkali feldspar and clinopyroxene increase towards higher H2Omelt (between 4 and 5 wt%). At 800 ºC major phase changes occur, with
co-crystallization of magnetite, clinopyroxene and biotite at liquidus conditions. Alkali feldspar is less affected and still crystallizes for a H2Omelt of 5 wt%, similar to
conditions at 850ºC. At 800ºC, ilmenite disappears and is replaced by titanite. The stability field of titanite is poorly constrained but the results show that under the T-fO2
conditions explored it only appears at >150 MPa (Figs. 4c and 4d).
Despite the small fO2 rangethat we have investigated we find significant changes
in phase stabilities with fO2. A decrease from NNO to NNO-0.5 at 850ºC and 100 MPa,
4.6 wt% at NNO-0.5) and moves the alkali feldspar field towards lower H2Omelt (Fig. 4,
Table 2).
5.2. Phase proportions
Calculated phase proportions show that, as expected, crystallinity increases with decreasing temperature, decreasing H2Omelt and increasing pressure (Fig. 4, Table 2). At
900ºC the crystal content is < 5 wt% for H2Omelt between 6 and 3 wt%, at all studied
pressures. For H2Omelt of about 1 to 3 wt% the crystal content increases to > 30 wt%. At
850ºC and H2Omelt in the range 6.5 - 5 wt% the crystal contents are < 5 wt%, but
crystallinities increase to up to 30 wt% when H2Omelt decreases to about 1 wt%
(between 200 and 100 MPa). The abrupt and strong increase in crystallinity is enhanced at 50 MPa, where small variations in H2Omelt (±0.5 wt%) increase the crystal content by
about 10 wt%. Charges run at 800ºC, 200 MPa, and close to water saturation conditions have crystal contents of 5 wt%, increasing up to 25 wt% when H2Omelt decreases by
only 0.5 wt%. Moreover, at such temperature and for a given H2Omelt, a decrease in
pressure from 200 to 100 MPa increases the crystal content by about 15 wt%. The large changes in crystallinities with varying H2Omelt are essentially controlled by the amounts
of alkali feldspar, whereas in charges with low crystal content (< 5 wt%) biotite, magnetite and ilmenite are the main crystallizing phases (Fig. 4, Table 2).
5.3. Mineral compositions
5.3.1. Ilmenite and magnetite
Ilmenite and magnetite were identified by SEM-EDS but they were often too small for reliable electron microprobe analysis to be performed. We nevertheless analysed Fe-Ti oxides in 7 charges, whose compositions have been corrected out from the glass contribution (less than 3 % of contamination) and the results are shown in Table 4. Magnetite FeO* and TiO2 contents range between 77-79 wt%, and17-14 wt%,
respectively, with an Mg# of 3-4. Ilmenite crystallizing at 850ºC and at H2Omelt 4-5 wt%
has TiO2 contents in the range 49-51 wt%, while its FeO* varies from 45 to 42 wt%,
depending on H2Omelt (Table 4).
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Experimental biotites have Mg# in the range 50-68, and define two populations depending on temperature (Fig. 5; Table 4): those from 900ºC have higher Mg# (> 57) than at 850ºC. At constant pressure and temperature, a decrease in H2Omelt decreases
Mg#. For example, at 900ºC and 50 MPa the Mg# decreases from 68 to 63 with a decrease in melt water content of only 0.7 wt%. Similar variations in Mg# are observed at 900ºC, 100MPa but, over a wider range of H2Omelt (Fig. 5). Otherwise, a decrease of
pressure at a constant temperature and H2Omelt increases the Mg# of biotite.
5.3.3. Clinopyroxene
Experimental clinopyroxenes are diopside and hedenbergite according to Morimoto (1989), spanning the range of compositions between En24 Fs26 Wo47 and En34
Fs19 Wo45, with Mg# varying between 47 and 64. Clinopyroxene composition is affected
by variations in experimental temperature and changes in melt water content (Fig.6). When temperature increases from 800 to 850ºC, the Mg# of clinopyroxene increases from 45 to 55, the En content from 24% to 29%, whilst Fs decreases from 27% to 23%, and Wo remains almost constant. These variations include charges ran at different P and fO2. At pressures 100 MPa a variation in H2Omelt does not appreciably
affect clinopyroxene composition, but in runs at 50 MPa, as H2Omelt increases, the En
and Mg# increase while Fs and Wo decrease (Fig. 6). Similar compositional variations were observed in experiments done on other phonolites (e.g., Andújar et al. 2008, 2010) but, their magnitude remain smaller compared to those observed in more silicic compositions (e.g., Scaillet and Evans, 1999).
5.3.4. Alkali feldspar
Alkali feldspars are anorthoclase with compositions between An9 Ab70 Or21 and
An2 Ab64 Or34, depending essentially on temperature and H2Omelt. At constant pressure,
both An and Ab decrease whereas Or increases with decreasing temperature and H2Omelt
(Fig. 7). However, the effect of temperature is larger than the effect of H2Omelt and the
compositional variations are enhanced at lower pressures. For example, at 50 MPa and for a constant H2Omelt of 2 wt%, an increase in temperature from 850ºC to 900ºC
increases the An content of anorthoclase from 2 mole % to 9 mole%. In contrast, at 200 MPa and constant H2Omelt (4 wt%), a variation in temperature from 850 to 800ºC does
5.4. Glass
Residual glass compositions are reported in Table 5. The Na2O+K2O is between
15 and 18 wt%, and SiO2 contents range between 61.7 and 58.3 wt% (reported
compositions were re-calculated to 100% anhydrous; Table 5; Fig. 8). According to Le Maitre et al. (1989), all experimental glasses are phonolitic and mildly per-alkaline (molar Na+K/Al between 1.1 and 1.3). The glass composition varies according to changes in intensive parameters and thus with phase proportions. For a given pressure and temperature, SiO2 displays a complex trend of increase and then decrease with
decreasing H2Omelt (Fig 8a). For a constant temperature, the highest SiO2 enrichment is
displaced towards higher H2Omelt with increasing pressure, though the effect is more
pronounced at < 900ºC (Fig. 8). The SiO2 variation is controlled by the onset of alkali
feldspar crystallization, although the maximum SiO2 enrichment depends also on the
relative abundances of oxides and biotite (Table 2). Glasses from charges at 200 MPa and 900ºC with biotite present have slightly lower SiO2 (60.5 wt%) than at 850ºC (61.5
wt%), where magnetite-ilmenite prevail over biotite (Figs.4 and 8). At a given H2Omelt,
the SiO2 content increases by about 0.5-1 wt% with an increase of 50ºC and of 50 MPa.
A decrease in temperature of 50ºC at a given pressure and H2Omelt, leads to a
decrease of TiO2, MgO and CaO contents by about 50% relative, whereas both Na2O
and K2O slightly increase. A decrease in pressure leads to an increase in MgO and a
decrease in CaO contents. However, at 50 MPa, compositional changes with intensive variables are different from those observed at higher pressures. Residual glasses at 900ºC have higher MgO contents than at 850ºC. Moreover, a decrease in H2Omelt
increases the MgO content of the glass. Such compositional variations have been also documented in other experiments performed on similar compositions at low pressure (e.g., Andújar and Scaillet, 2012a). The effect of fO2 on glass compositions is small,
reflecting the limited range of fO2 investigated. However, experiments conducted at
more reducing conditions (NNO-2) yield glasses with lower TiO2, MgO and CaO and
higher SiO2, Al2O3 and K2O contents than those synthesised at NNO.
6. Discussion
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The comparison between the mineral assemblage of the natural phonolite (14 wt % of phenocrysts, mainly alkali feldspar, but also clinopyroxene, biotite and Fe-Ti oxides) and those from experiments provides first-order constraints on the magma pre-eruptive conditions. At temperatures between 850 and 900ºC the crystal content and phase assemblage of the natural phonolite have been reproduced at pressures between 75 and 25 MPa, and H2Omelt 2.0-4.0 wt% (Fig. 4). Lower temperatures (e.g. 800 ºC) can
be ruled out because of the much higher crystal content and absence of ilmenite in the explored ranges of H2Omelt and pressures. Pressures higher than about 75 MPa can also
be ruled out because clinopyroxene only occurs at crystal contents higher than in the natural phonolite. We have not performed experiments at < 50 MPa, but extrapolation of the crystallinity and phase equilibrium relationships obtained at 50 MPa also suggests much higher crystallinities at pressures significantly lower than 50 MPa. To refine and test the robustness of such constraints we use below the compositions of mineral and glass.
Natural magnetite has a Mg# of 5 whereas our two experimental magnetites from experiments conducted at 850ºC, 100 and 50 MPa, and H2Omelt of about 3 wt%
have Mg# of 3 to 4. Magnetite composition varies with temperature and fO2 (e.g,
Andújar et al. 2008, 2010) and thus, the slightly higher Mg# of the natural magnetite suggests higher temperatures (e.g., > 850ºC) or/and more oxidizing conditions (e.g., fO2
> NNO). We have reproduced the TiO2 content of the natural ilmenite at 850ºC, 200
MPa and H2Omelt of 5.5 wt%, although these conditions are different from those inferred
above.
Experimental biotites crystallizing between 50 and 200 MPa at 900ºC, and H2Omelt between 1.5-4.5 wt% encompass the Mg# of natural biotites (Mg# 63, Fig. 5).
However, at this temperature only experiments performed at 50 MPa and H2Omelt < 2.5
wt% reproduce the phase assemblage and crystal content of the natural phonolite. At 850ºC the experimental biotites have lower Mg# than those from the phonolite.
We have successfully reproduced the rim and core compositions of the natural alkali feldspars. Their Ca-rich rims (An7-8)are reproduced at 900ºC, 50 MPa and H2Omelt 2 wt% while cores (the bulk of crystals, see Fig 3), which range from An2.5 to An4, are
reproduced at 850ºC, 50 MPa and H2Omelt between 2-3 wt% (Fig. 7). These results
at 900ºC. Such a temperature difference suggests a change in the storage temperature of the reservoir just prior to eruption which, as we show below, is also consistent with biotite and Fe-Ti oxide reequilibration.
We have not obtained EMP analyses of clinopyroxene in charges produced at 900ºC due to the small size of crystals; however, the compositional trends displayed by the experimental clinopyroxenes support this temperature for natural clinopyroxene crystallization. At 850ºC, 50 MPa and H2Omelt of 2 wt%, both the Mg# and En content
of natural clinopyroxene are closely reproduced yet, Fs and Wo are slightly different. However, an increase of temperature of 50ºC will produce a 5% moles increase in the En content and Mg# of clinopyroxenes but also, a 5% moles decrease in the Fs content, producing a clinopyroxene similar to the natural one (Fig.6; see previous section).
The comparison between natural and experimental phase relationships and mineral compositions therefore suggests pre-eruptive conditions for the Roques Blancos phonolite to be 850-900ºC, 50 MPa, and melt water contents between 2 to 2.5 wt% (Ptotal< PH2O). At such conditions the SiO2, CaO, Na2O, MgO, TiO2 and K2O contents
of the residual melt are closely reproduced (Table 1; Fig. 8). Only FeO* and Al2O3
contents are different (but almost within error if analytical uncertainties are considered) than those from experiments conducted at 850-900ºC and 50 MPa (Fig. 8).
We do not have enough experimental data from different fO2 (i.e., biotite, Fe-Ti
oxides) to make a precise determination of the pre-eruptive oxygen fugacity. However, the fact that we have successfully reproduced the phase assemblage, crystal content and phase compositions of the natural phonolite at an fO2NNO suggests that such an fO2
broadly prevailed in the reservoir. Moreover, the Mg# of the natural biotite (which is highly sensitive to fO2) is reproduced at an fO2NNO-1; hence a pre-eruptive fO2
between NNO and NNO-1 appears to be a good redox estimate for the natural phonolite, in accord with constraints obtained from natural magnetite-ilmenite pairs (Table 1) and with those obtained previously for other phonolites from Tenerife (Ablay et al., 1998; Andújar et al., 2008, 2010; Andújar and Scaillet, 2012a). On the basis of the foregoing discussion we conclude that prior to eruption the Roques Blancos phonolite was stored at 50±15 MPa, 875oC ± 25 oC , with H2Omelt 2-2.5 wt%, at an fO2 of
NNO-0.5 (±NNO-0.5). Although the pre-eruptive H2Omelt is rather low, the low pressure of magma
17
shows that, if present, CO2 must reside primarily in the vapor phase, albeit in moderate
amounts to be compatible with our inferred H2Omelt.
6.2. Evidence for late magma reheating upon mixing before eruption
Many Roques Blancos alkali feldspars are characterised by a reverse compositional zoning with An-rich rims and Or-poor cores (Fig. 3). We have successfully reproduced such An-rich compositions at 900ºC, 50 MPa, 2 wt% H2Omelt
whereas cores are reproduced at the same pressure and H2Omelt but at a lower T, 850°C
(Figs. 4 and 7). Reverse zoning in feldspars are generally interpreted as arising from crystallisation of a more Ca-rich melt, often related to mixing with a more mafic magma (e.g., Couch et al. 2001, Martel et al, 2006). Alternatively, experimental works have shown that such Ca-rich rims can be also produced by increasing temperature without interaction with a mafic body (e.g., Rutherford et al., 1985; Couch et al., 2001), by increasing pressure at water-saturated conditions (Rutherford and Devine 2008), or by an increase in water activity via the preferential loss of a CO2 (Holloway, 1976). To
discriminate between the different possibilities, it is useful to compare changes in An content with changes in other minor elements as shown by Triebold et al. (2006) and Ruprecht and Wörner (2007). These authors suggest that a process of magma mixing can be identified in feldspars when the An enrichment is accompanied by an increase in Fe content whereas when Fe remains constant, only an increase in temperature can explain the Ca-rich compositions. The compositional traverses show that the increase in the An content towards the rims of the crystals is indeed coupled with an increase in the FeO*, TiO2 and BaO which suggests that such Ca-rich rims record the income of a more
mafic and higher temperature melt, as proposed by Triebold et al. (2006) and Ruprecht and Wörner (2007). This hypothesis is also supported by our experimental results: the observed CaO-FeO*-TiO2 enrichment of the alkali feldspar rims cannot be explained
only by a single shift in temperature or water activity because in each case, the stability field, phase proportions and composition of phases like biotite, magnetite, or clinopyroxene will be also modified. In detail, at 900ºC for reproducing the low Ca-cores of feldspars, crystal contents > 30 wt% and H2Omelt <1.7 wt% are required (fig. 4).
Then, to generate the An7-8 rims, H2Omelt should increase to 2.5 wt% and the crystal
H2Omelt will increase the Ca content of the coexisting melt, producing a shift in the An
content, it will keep FeO* and TiO2 concentrations constant, a trend unlike the rimward
enrichment in these elements observed in natural feldspars (Fig. 8). The above lines of evidence thus suggest that the most likely mechanism for explaining the compositional zonation of alkali feldspars is mixing with a hotter and more mafic magma. In such a scenario, the observation that only biotite and Fe-Ti oxides compositions record the conditions following the intrusion of the more mafic and hotter melt can be explained by their much faster kinetics of intracrystalline diffusion compared to other minerals (e.g., Cherniak and Dimanov, 2010; Van Orman and Crispin, 2010). Although we do not have detailed compositional zoning in clinopyroxene to further check this hypothesis, the BSE images show multiple growth zones with dissolution surfaces, which is consistent with an open-system evolution with abrupt changes in temperature and composition.
Magma mixing and mingling are common processes occurring at Tenerife, as recorded by the existence of mingled-mixed products or by reversely zoned phenocrysts (Araña et al., 1994; Ablay et al., 1998; Neumann et al., 1999, Triebold et al, 2006), as discussed above. The steep increase in Ba recorded in the alkali feldspars (Fig. 3) provides insights into the composition of the mafic end-member. According to Ablay et al. (1998), among the mafic magmas erupted in the last 30 kyr, tephriphonolites are the most Ba-enriched, in particular those erupted from Pico Viejo vent, which can contain up to 1400 ppm of Ba. Estimated pre-eruptive temperatures of such tephriphonolites are in the range 1020-1040°C (Ablay et al., 1998; Neumann et al., 1999), hence they are hot enough to increase temperature, at least locally, yielding Ca-rich overgrowth on alkali feldspars of the Roques Blancos phonolite when entering in the reservoir, although in detail the thermal effects will depend on the relative masses of juxtaposed magmas.
19
more mafic intrusion might have triggered the eruption some time later (e.g., Folch and Martí, 1998).
6.3. Central versus flank eruptions in Tenerife and their relation to different eruptive styles
The volcanic activity of the last 30 kyr of the Teide-Pico Viejo complex is characterised by the eruption of intermediate to phonolitic magmas which volumetrically predominated over the basaltic ones (Ablay et al., 1998; Carracedo et al., 2007, Martí et al., 2008). The phonolitic eruptions alternate between central Pico Teide and Pico Viejo vents and several flank domes located around the stratovolcanoes, with these last dominating the evolution of the actual volcanic system in terms of number of events (Ablay and Martí, 2000; Carracedo and Badiola, 2006; Rodríguez-Badiola et al., 2006; Carracedo et al., 2007). Flank dome eruptions share several physical, petrographic, geochemical, and volcanological features. Phonolitic domes are located at heights between 2000 and 3000 m a.s.l. over the steep flanks of the Teide-Pico Viejo stratovolcanoes: generally they are volumetrically smaller than eruptions occurring from the central vents (Ablay, 1997; Martí et al., 2008; Carracedo et al., 2007; Martí et al., 2012) and are characterised by the emission of thick phonolitic lava flows. They typically have also an explosive phase that produces significant fall-out deposits, as documented for the eruptions of Montaña Blanca and El Boquerón (Ablay et al. 1995; Andújar and Scaillet, 2012a; García et al., 2012)
Previous estimates of the storage conditions of flank dome eruptions in Tenerife were done using thermo-barometric models (Ablay, 1997; Ablay et al., 1995; Ablay et al., 1998). Coexisting Fe-Ti oxides of Arenas Blancas, Montaña Las Lajas, and Montaña Blanca flank eruptions yield temperatures in the range 775-900°C (Ablay, 1997; Ablay et al., 1995; Ablay et al., 1998). These authors also measured water contents of melt inclusions trapped in clinopyroxenes from Montaña Majua and Montaña Blanca flank dome eruptions, reporting values of 1.2 - 2.5 wt%, and 3 - 4.5 wt%, respectively. Using the volatile content and the water solubility model of Carroll and Blank (1997), Ablay et al. (1995) inferred a storage pressure of 100 MPa for the Montaña Blanca eruption. New experimental results by Andújar and Scaillet (2012a) using the products of the sub-plinian phase of the Montaña Blanca eruption suggest instead that this magma was stored at 850±15°C, 50±15 MPa, 2.5±0.5 wt% H2Omelt. This phase of the eruption
with a composition similar to that of Roques Blancos. We have calculated the magma viscosities of these two phonolites (models of Giordano et al., 2008, and Scaillet et al. 1998; Andújar and Scaillet, 2012b) and have found about the same values (104.1 Pa*s). This and their similar storage conditions, which are close to water-saturation according to available H2O-solubility models in phonolitic melts (Carroll and Blank, 1997;
Schmidt and Behrens, 2008), indicate that both magmas would respond similarly upon decompression. The explosive nature of the Montaña Blanca eruption suggests that it is likely that the beginning of the Roques Blancos eruption had a short-lived highly explosive initial phase, which is in agreement with the recent identification of fall-out deposits associated with this eruption (O. García unpublished data).
The physical and petrological features shared between the most recent satellite phonolitic eruptions of Teide-Pico Viejo (Fig. 1) and that of Montaña Blanca, El Boquerón and Roques Blancos, all point towards a shallow origin for these flank eruptions, which appear to have been fed by reservoirs located at depths of 1-2 km below the surface. These shallow values contrast with those experimentally determined for central eruptions occurring from Teide proper, which were constrained to be about 5 km below the actual summit of Pico Teide volcano (Andújar et al., 2010). In figure 9 we have projected the depths for the Roques Blancos, Montaña Blanca and central Teide phonolitic magmas. It shows that phonolitic magmas can be stored at various depths in Tenerife, up to very close to the surface (~1-2 km below the surface), suggesting that the phonolite storage region of Teide is not characterised by one single and large reservoir, but instead by various isolated, possibly transient, magma pockets that can coexist simultaneously, re-enforcing and supporting previous observations obtained from petrological data and modelling efforts (Ablay et al., 1998; Andújar and Scaillet, 2012a; Martí and Geyer, 2009). The characteristics of the actual plumbing system beneath Teide-Pico Viejo where several phonolitic shallow reservoirs can co-exist , as well as the difference in storage depths, can thus explain the multiple shifts between central and lateral eruptions documented to occur at this stratovolcano (Martí and Geyer, 2009).
7. Conclusions
21
assemblage, phase compositions) of the natural phonolite are reproduced at 900°C, though the composition of Ca-poor cores of alkali feldspar record lower temperatures of about 850°C. The zoning pattern of natural alkali feldspars show that they have high Ba, Ti, and Fe rims which, we suggest, record a late and transient reheating following mixing of the resident magma with a more mafic melt. Such shallow and water-rich storage conditions were also determined for another flank and Plinian eruption of Teide and contrast with the deeper reservoir inferred for predominantly effusive eruptions of magmas emitted from the central Teide cone. Such differences in storage depths and locations illustrate the complexity of the magma plumbing system that may currently exist below Teide, which does not appear to be fed from a single large reservoir. Instead, the evidence call for multiple isolated small magma reservoirs, whose level of assembly in upper crust is in part controlled by the feedback existing between the evolution and growth of the aerial cone built/destroyed during eruptions and the local stress field at depth.
The results from the current work should help improve hazard assessment at Tenerife, along with the forecasting the future behaviour of phonolitic magmas stored below Teide-Pico Viejo volcanic system. In particular, the new constraints we provide about both pre-eruptive H2Omelt of phonolites and their storage depths may help in
providing a first estimate of the explosive/effusive potential of these magmas by using the location of tremors during an on-going eruption at central parts of the island. Nevertheless, we stress that the explosive vs non-explosive character of alkali-rich magmas can be easily changed by additional factors (ie, influx of other volatile species like CO2, degassing, interaction with meteoric water, mixing with mafic compositions
within the conduit); and will be always subordinated to the composition of the ascending magma, which is so far virtually impossible to know with the available monitoring tools.
Acknowledgements
logistical support during rock sampling at Tenerife, as well as for the comments and editorial handling of the paper.
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Figure 4. Isothermal phase relations of the Roques Blancos phonolite at a) 900ºC and
fO2NNO; b) 850ºC and fO2NNO; c) 800ºC and fO2NNO for different pressures and
water contents in the melt; d) Isobaric phase relations (50 MPa and fO2NNO) for various temperatures and water contents in the melt. Mag: magnetite, Bt: biotite, Cpx: clinopyroxene, ilm: ilmenite, tit: titanite, fsp: alkali feldspar. Dashed lines are estimated phase boundaries. Numbers above dashed lines indicate crystal content in wt (%). Label
An below squares in plate d indicates the An content of experimental feldspars in the
Figure 5. a) Variation of the Mg# with water content in the melt of natural and experimental biotites. Black line and grey horizontal bar shows the natural composition. Numbers next to symbols in the legend indicate temperature, pressure and fO2
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Figure 8. Glass compositional variations of major and minor oxides with water content in the melt (plates a to h). Black line and grey horizontal bar shows the natural composition. Numbers next to symbols in the legend indicate temperature, pressure and
Table 1: Major element compositions (in wt%) of bulk-rock, glass, and minerals of the natural sample, and starting material.
bulk rock a starting material b
calculated
glass c sd magnetite sd anorthoclase sd diopside sd biotite sd ilmenite sd
n 15 6 17 8 10 6 SiO2 59,38 59,64 59,23 0,33 0,08 0,01 65,11 0,40 51,36 0,41 36,66 0,28 0,05 0,01 TiO2 0,66 0,74 0,63 0,06 16,71 0,58 0,09 0,02 0,85 0,13 7,31 0,15 48,45 0,22 Al2O3 18,89 19,41 19,01 0,23 1,00 0,01 19,72 0,29 1,37 0,18 12,55 0,15 0,10 0,00 MgO 0,41 0,48 0,41 0,05 2,09 0,09 0,00 0,00 12,59 0,39 13,87 0,17 2,96 0,12 CaO 0,88 0,80 0,86 0,09 0,16 0,20 0,93 0,20 21,66 0,24 0,00 0,00 0,10 0,04 MnO 0,19 0,14 0,21 0,02 2,50 0,01 0,00 - 0,83 0,03 0,00 0,00 3,26 0,03 FeO* 3,41 3,39 3,62 0,18 72,13 0,23 0,25 0,03 10,07 0,52 14,60 0,19 42,95 0,55 Na2O 9,90 9,90 10,37 0,19 - - 7,94 0,19 1,16 0,08 1,07 0,12 - -K2O 5,48 5,49 5,65 0,12 - - 4,81 0,42 0,00 0,00 8,45 0,08 - P2O5 0,07 - - - -Cl 0,30 - - - -F 0,12 - - - -Original Sum 99,69 100,0 100,0 94,7 0,2 98,9 0,3 99,9 0,6 94,5 97,9 0,0 Mg# 16,96 4,90 69,03 1,01 62,87 0,39 An 4,44 0,95 Ab 68,30 1,47 Or 27,26 2,35 Wo 49,6 0,3 En 40,1 1,1 Fs 10,3 1,2 Phase proportions (wt%) 85,60 0,30 13,70 0,08 0,30 0,10
Thermo-Oxybarometric results after Sauerzapf et al. (2008)
n. pairs 6 sd
T(ºC) 895,37 4,8
∆NNOd -0,19 0,10
log f O2 -12,19 0,02
a: Bulk rock composition of the natural sample analysed by ICP-MS. b: Electron Microprobe analysis of the starting glass added to the capsules.
c: Composition of the residual melt calculated by mass balance using the composition and abundance of the mineral phases and the bulk rock. sd: Standard deviation.
n: Number of analysis. *: Total iron reported as Fe2+. Mg# =100[Mg/(Mg+Fe*)].
Or, Ab, and An end- members calculated as in Deer et al. (1972)
End members for clinopyroxene (En, Fs, Wo) calculated as in Morimoto (1989). phase proportions calculated as weight %.
Table 2: Experimental conditions and run products
NNO experiments
Run XH2Owt%in H2Owt%melt f H2(bar) log f O2(bar) D NNO aH2O
900ºC,200MPa,164h a 8,36c -11,69 0,27 RB85 1,00 5,74 -12,11 -0,19 0,69 (97.5)gl+(2.5)bt 0,73 RB86 0,93 5,33 -12,23 -0,31 0,60 (97.5)gl+(2.5)bt 0,43 RB87 0,83 4,74 -12,44 -0,52 0,47 (97.3)gl+(2.7)bt 0,39 RB88 0,70 4,00 -12,74 -0,81 0,34 (97.8)gl+(0.9)mag+(trace)ilm+(1.3)bt 0,06 RB89 0,58 3,35 -13,04 -1,12 0,24 (97.7)gl+(1.4)mag+(trace)ilm+(0.4)bt 0,35 RB90 0,44 2,53 -13,52 -1,60 0,14 (88.1)gl+(0.9)mag+(trace)ilm+(0.7)bt+(10.3)fsp 0,07 900ºC,100MPa,188h a 4,89d -11,81 0,15 RB 91 1,00 4,53 -11,94 0,00 0,86 (98.4)gl+(1.6)bt 0,34 RB 93 0,82 3,73 -12,27 -0,33 0,59 (98.2)gl+(1.5)mag+(trace)ilm+(0.3)bt 0,37 RB 94 0,73 3,31 -12,48 -0,54 0,46 (98.2)gl+(0.6)mag+(0.2)ilm+(0.7)bt 0,30 RB 95 0,56 2,53 -12,94 -1,00 0,27 (89.8)gl+(1.0)mag+(trace)ilm+(0.3)bt+(8.9)fsp 0,11 900ºC,50MPa,284h a 3,53c -11,75 0,21 RB 96 1,00 3,07 -11,95 0,01 0,80 (98.2)gl+(1.8)bt 0,30 RB 97 0,91 2,78 -12,11 -0,16 0,66 (98.4)gl+(1.6)bt 0,57 RB 98 0,78 2,39 -12,37 -0,41 0,49 (98.5)gl+(0.7)mag+(trace)ilm+(0.8)bt 0,26 RB 99 0,73 2,24 -12,48 -0,52 0,43 (85.0)gl+(1.3)mag+(0.1)ilm+(0.1)bt+(trace)cpx+(13.5)fsp 0,08 RB 100 0,60 1,85 -12,79 -0,84 0,30 (74.3)gl+(1.3)mag+(trace)ilm+(0.4)bt+(0.2)cpx+(23.8)fsp 0,10 RB 101 0,50 1,53 -13,09 -1,13 0,21 (67.8)gl+(1.8)mag+(trace)ilm+(0.6)bt+(0.2)cpx+(29.6)fsp 0,09 QFM experiments 850ºC,100MPa,188h b 6,66c -13,03 -0,14 RB 73 1,00 5,09 -13,25 -0,38 1,00 (98.0)gl+(0.7)mag+(0.5)ilm+(0.8)bt 0,70 RB 74 0,90 4,57 -13,32 -0,45 0,89 (98.3)gl+(0.4)mag+(0.7)ilm+(trace)bt+(0.6)cpx 0,58 RB 75 0,80 4,09 -13,48 -0,61 0,67 (88.7)gl+(0.9)mag+(trace)ilm+(0.9)bt+(0.5)cpx+(9.0)fsp 0,25 RB 76 0,71 3,61 -13,73 -0,86 0,51 gl+mag+ilm+bt+cpx+fsp n.d. RB 77 0,55 2,80 -13,88 -1,01 0,31 gl+mag+ilm+bt+cpx+fsp n.d. RB 78 0,53 2,70 -14,15 -1,28 0,26 gl+mag+ilm+bt+cpx+fsp n.d.
XH2Owt%in:initial H2O/(H2O+CO2) in the charge
H2Owt%melt: water content in the melt a: determined by FTIR, b:determined by the solubility model of Papale et al. (2006)
f H2(bar):hidrogen fugacity of the experiment. c:determined by using NiPd alloy sensors, d:calculated using the data obtained from succesfull NiPd alloys (see text for details) log f O2(bar): logarithm of the oxygen fugacity calculated from the experimental f H2
D NNO:log f O2-log f O2 of the NNO buffer calculated at P and T (Pownceby & O'Neill,1994) e:aH2Ocalculated by using H2Owt% at saturation/ H2Owt% in the melt (see text for details) n.d.: not determined
Table 3: FTIR data of experimental runs
Sample n Thickness(µm) sd 5200(cm-1) sd 4500(cm-1) sd Density(g/l)a sd H2O(wt%)b sd
RB85 3 197,0 2,1 0,05 0,00 0,01 0,00 2579,5 12,0 5,72 0,60
RB91 3 135,0 3,5 0,07 0,00 0,03 0,00 2641,5 2,6 4,53 0,12
RB96 3 91,5 2,6 0,03 0,00 0,02 0,00 2672,0 5,3 3,07 0,23
n: number of spectra sd:standard deviation
5200(cm-1): Absorvance intensity from the 5200 peak 4500(cm-1):Absorvance intensity from the 4500 peak a: density of the melt calculated as in
Table 4: Composition (wt%) and end-members of experimental mineral phases
Magnetite
SiO2 TiO2 Al2O3 MgO CaO MnO FeO Na2O K2O Sum Mg# wt%gl
RB32(3) - 17,44 0,91 1,54 0,00 3,02 77,13 - - 100,0 3,43 1,20
sd - 0,29 0,01 0,09 0,00 0,10 0,26 - - 0,0
-RB16 - 13,75 2,10 1,75 0,00 3,53 78,87 - - 100,0 3,81 2,10
Ilmenite
SiO2 TiO2 Al2O3 MgO CaO MnO FeO Na2O K2O Sum Mg# wt%gl
RB13(3) - 49,49 0,10 2,22 0,00 3,75 44,47 - - 100,0 - 0,50 sd - 0,15 0,05 0,15 0,00 0,10 0,13 - - 0,0 -RB14(4) - 50,10 0,12 2,31 0,00 4,19 43,28 - - 100,0 - 2,30 sd - 0,42 0,06 0,06 0,00 0,08 0,48 - - 0,0 -RB15 - 50,47 0,07 2,43 0,00 4,86 42,22 - - 100,0 - 3,00 RB3 - 48,72 0,07 1,98 0,00 3,47 45,72 - - 100,0 - 1,10 RB77(2) 0,0 0,00 51,06 0,00 1,78 4,56 42,85 - - 100,0 -Biotite
SiO2 TiO2 Al2O3 MgO CaO MnO FeO Na2O K2O Sum Mg# wt%gl
-RB89(2) 36,6 5,58 16,68 14,17 0,36 11,17 0,07 1,41 8,07 95,1 58,43 -sd 0,5 0,11 0,67 0,57 0,06 0,25 0,08 0,05 0,02 0,8 0,43 -RB94 38,2 6,18 14,59 14,51 0,47 12,16 0,03 1,34 8,16 96,6 59,90 -RB99(2) 36,2 5,54 15,54 12,34 0,41 14,70 0,04 1,22 8,28 95,2 67,98 -sd 0,5 0,41 0,22 0,44 0,13 0,22 0,01 0,14 0,21 0,6 0,45 -RB100 (2) 37,5 6,39 14,71 12,45 0,38 13,66 0,02 1,22 8,53 96,4 66,17 -sd 0,3 0,14 0,90 0,22 0,14 0,11 0,02 0,20 0,13 0,5 0,58 -RB101 39,0 5,71 15,01 13,31 0,52 12,85 0,08 2,13 8,40 97,0 63,25 -Clinopyroxene
SiO2 TiO2 Al2O3 MgO CaO MnO FeO Na2O K2O Sum Mg# wt%gl En Fs Wo
RB32(4) 48,4 2,45 3,15 9,36 20,49 1,00 12,27 1,62 0,09 98,8 57,62 - 29,7 21,8 46,7 sd 0,3 0,22 0,21 0,20 0,14 0,06 0,14 0,13 0,02 0,3 0,59 - 0,5 0,2 0,3 RB34(4) 49,9 1,60 1,85 11,43 21,00 0,97 11,58 1,66 0,00 100,0 63,75 - 34,0 19,4 44,9 sd 0,3 0,30 0,83 0,32 0,80 0,11 0,21 0,32 0,00 0,0 0,36 - 0,3 0,4 0,9 RB49(3) 50,3 0,79 1,30 7,95 21,87 1,48 15,42 0,93 0,00 100,0 47,89 2,80 24,0 26,1 47,4 sd 0,6 0,30 0,32 0,36 0,07 0,12 0,45 0,07 0,00 0,0 1,85 1,73 1,0 0,9 0,3 RB52 50,4 0,67 1,18 7,59 20,96 1,49 15,23 1,38 0,14 99,0 47,04 - 23,7 26,7 47,0 RB76(2) 51,0 1,58 4,04 8,65 19,60 1,17 12,17 1,91 0,47 100,5 55,89 - 28,6 22,6 46,6 sd 0,4 0,48 0,88 0,11 0,47 0,05 0,45 0,18 0,08 1,3 1,24 - 0,8 0,5 0,4 RB77(2) 50,1 0,96 1,93 8,75 20,64 1,29 12,92 1,41 0,19 98,2 54,68 - 27,7 23,0 47,0 sd 0,4 0,12 0,60 0,40 0,30 0,05 0,25 0,25 0,02 0,1 1,61 - 0,8 0,8 0,1 Feldspars
SiO2 TiO2 Al2O3 MgO CaO MnO FeO Na2O K2O Sum Mg# wt%gl An Ab Or
RB4(2) 64,1 0,36 19,64 0,18 1,23 0,00 1,50 7,75 4,14 98,9 - - 6,1 69,5 24,4
sd 0,7 0,14 0,48 0,02 0,04 0,00 0,54 0,20 0,09 0,8 - - 0,0 0,0 0,1
RB5 64,4 0,44 19,24 0,17 0,94 0,00 1,04 7,76 4,90 98,9 - - 4,5 67,5 28,0
RB31 65,9 0,20 19,64 0,03 1,08 0,02 0,53 7,72 5,03 100,1 - - 5,12 66,41 28,5
RB51(3) 64,6 0,16 20,24 0,03 0,95 0,00 0,44 7,63 4,38 98,4 - - 4,7 69,1 26,1 sd 0,1 0,01 0,36 0,02 0,01 0,00 0,22 0,06 0,07 0,5 - - 0,0 0,4 0,5 RB99 64,6 0,16 20,89 0,03 1,92 0,00 0,12 8,40 3,83 100,1 - - 8,8 70,1 21,1 RB100 64,3 0,23 20,78 0,00 1,66 0,13 0,52 8,26 4,14 100,1 - - 7,7 69,4 22,9 RB101(2) 64,1 0,28 19,60 0,07 0,78 0,19 0,73 7,54 5,53 99,3 - - 3,69 64,95 31,4 0,7 0,02 0,12 0,00 0,04 0,04 0,16 0,22 0,06 0,6 - - 0,12 0,29 0,41
numbers in brackets indicate the number of analysis. sd: Standard deviation
* Total Iron reported as FeO Mg# = 100[Mg/(Mg+Fe*)]
wt% glass: Clinopyroxene compositions are re-calculated due to glass contamination (see text for details) and here we indiciate the wt% glass substracted to the original electron microprobe analysis
En, Fs, Wo: calculated as in Morimoto (1989)
Table 5: Experimental melt composition (wt%) normalized to 100% anhydrous basis Run
NNO experiments
850ºC,200MPa,162h
SiO2 TiO2 Al2O3 MgO CaO MnO FeO* Na2O K2O Sum Mg#
original
sum Na2O+K2O per-alk
RB 100(5) 58,8 0,69 19,4 0,46 0,79 0,18 3,00 10,5 5,8 100 21,5 97,8 16,4 1,2 sd 0,4 0,05 0,2 0,06 0,05 0,09 0,30 0,2 0,1 - 1,3 RB 101(5) 58,7 0,70 19,6 0,46 0,78 0,17 2,60 10,9 5,9 100 24,0 98,4 16,8 1,2 sd 0,4 0,06 0,2 0,05 0,04 0,07 0,11 0,3 0,2 - 0,7 QFM experiments 850ºC,100MPa,188h RB 73(3) 60,8 0,48 19,9 0,18 0,92 0,17 2,61 9,4 5,6 100 10,8 93,7 15,0 1,1 sd 0,3 0,03 0,0 0,03 0,06 0,15 0,09 0,1 0,2 - 0,9 RB 74(3) 60,5 0,41 20,0 0,21 0,79 0,24 2,78 9,4 5,7 100 11,9 93,9 15,1 1,1 sd 0,3 0,03 0,0 0,03 0,06 0,15 0,09 0,1 0,2 - 0,8 RB 75(3) 59,8 0,50 20,0 0,19 0,88 0,23 2,62 10,2 5,6 100 11,5 95,0 15,7 1,1 sd 0,3 0,03 0,0 0,03 0,06 0,15 0,09 0,1 0,2 - 0,9
numbers in brackets indicate the number of analysis
All analyses are normalized to 100% anhydrous. Original totals are reported (original sum) sd: standard deviation
Mg#: magnesium number of the melt calculated as Mg#=100*(Mg/Mg+Fe) in moles per-alk: per-alkalinity calculated as (Na+K)/Al in moles
